Hydraulic conductivity of fractured upper crust: insights from hydraulic tests in boreholes and fluid-rock interaction in crystalline basement rocks

$Page 1: Hydraulic conductivity of fractured upper crust: insights from hydraulic tests in boreholes and fluid-rock interaction in crystalline basement rocks$
Hydraulic conductivity of fractured upper crust: insightsfrom hydraulic tests in boreholes and fluid-rock interactionin crystalline basement rocks

I . STOBER1 AND K. BUCHER2

1Karlsruhe Institute of Technology KIT, Institute of Applied Geosciences, Karlsruhe, Germany; 2Mineralogy and Petrology,

University of Freiburg, Freiburg, Germany

ABSTRACT

The permeability (j[m2]) of fractured crystalline basement of the upper continental crust is an intrinsic property

of a complex system of rocks and fractures that characterizes the flow properties of a representative volume of

that system. Permeability decreases with depth. Permeability can be derived from hydraulic well test data in deep

boreholes. Only a handful of such deep wells exist on a worldwide basis. Consequently, few data from hydrauli-

cally tested wells in crystalline basement are available to the depth of 4–5 km. The permeability of upper crust

varies over a very large range depending on the predominant rock type at the studied site and the geological his-

tory of the drilled crystalline basement. Hydraulic tests in deep boreholes in the continental crystalline basement

revealed permeability (j) values ranging over nine log-units from 10�21 to 10�12 m2. This large variance also

decreases with depth, and at 4 km depth, a characteristic value for the permeability j is 10�15 m2. The perme-

ability varies with time due to deformation-related changes of fracture aperture and fracture geometry and as a

result of chemical reaction of flowing fluids with the solids exposed along the fractures. Dissolution and precipita-

tion of minerals contribute to the variation of the permeability with time. The time dependence of j is difficult to

measure directly, and it has not been observed in hydraulic well tests. At depths below the deepest wells down

to the brittle ductile transition zone, evidence of permeability variation with time can be found in surface expo-

sures of rocks originally from this depth. Exposed hydrothermal reaction veins are very common in continental

crustal rocks and witness fossil permeability and its variation with time. The transient evolution of permeability

can be predicted from models using fictive and simple starting conditions. However, a geologically meaningful

quantitative description of permeability variation with time in the deeper parts of the brittle continental crust

resulting from combined fracturing and chemical reaction appears very difficult.

Key words: brittle–ductile transition, deep well, hydrothermal, permeability, veins, well test

Received 13 January 2014; accepted 20 August 2014

Corresponding author: Ingrid Stober, Karlsruhe Institute of Technology KIT, Institute of Applied Geosciences,

Adenauerring 20b, Karlsruhe D-76131, Germany.

Email: [email protected]. Tel: +49 721 608 45488.

Geofluids (2015) 15, 161–178

INTRODUCTION

Features of brittle deformation, such as fractures, faults,

joints, and veins, are the principal water (fluid)-conduct-

ing structures in crystalline basement rocks and provide

the dominant conduits for fluid flow in the brittle upper

continental crust. Fluid flow in this environment can be

described by the flow law of Darcy. Flow is driven by a

hydraulic gradient and controlled by the prime parame-

ters of advective fluid, solute, and heat transport in frac-

tured crystalline rocks, namely permeability and porosity.

Depending on the scale of interest, different methods

are used to determine these two parameters. These

include laboratory measurements (e.g., Berckhemer et al.

1997), fracture analysis (Caine et al. 1996; CFCFF

1996; Jakob et al. 2003; Mazurek et al. 2003; Konzuk

& Kueper 2004), geophysical modeling of heat flow data

(e.g., Hayba & Ingebritsen 1994), and in situ testing of

boreholes (e.g., Cooper & Jacob 1946; Ferris 1951;

Butler 1998; Nielsen 2007; Peters 2012). Well tests pro-

vided permeability data from the crystalline basement to

depths of 5 km and offer insights into the permeability

© 2014 John Wiley & Sons Ltd

Geofluids (2015) 15, 161–178 doi: 10.1111/gfl.12104

structure of the crust and its variation with depth and

lithology.

It is important, however, to critically analyze and evalu-

ate permeability data derived from well tests reported in

the literature before they can be used (or discarded) in

large-scale models of the permeability structure of the

crust. In this study, we address several aspects and diffi-

culties in deriving permeability data from hydraulic well

tests.

Porosity and permeability are subject to unceasing

changes in the brittle upper continental crust. Reactive

fluid flow along the fractures continuously modifies the

aperture and surface structure of the fractures mostly by

chemical dissolution and precipitation reactions (Weisen-

berger & Bucher 2011; Bucher et al. 2012; Alt-Epping

et al. 2013). After initial porosity and permeability genera-

tion by creation of a fracture, the permeability tends to

increase during a certain period. Later, permeability and

porosity decreases due to precipitation of secondary miner-

als along the fractures until they become completely sealed.

This fracture–reaction–seal process of reaction vein forma-

tion may occur in several cycles and cause incremental vein

growth. Although there are ample field observations that

imply the significance of the described permeability–time

relationship in fractured upper crust (e.g., Bucher-Nurmi-

nen 1981), the time-dependent permeability evolution can-

not normally be recorded by hydraulic well tests because of

incompatible timescales of the fracture–reaction–seal pro-

cess (Bucher-Nurminen 1989) and the well test methods.

Nevertheless, we try to illustrate in this study structures

characteristic of transient permeability in fractured upper

crustal rocks.

Fluid flow in fractured upper crustal aquifers is generally

driven by hydraulic gradients, which may result from a

number of different feasible causes and imbalances includ-

ing topography, thermal, and chemical disequilibrium.

Pumping or injection tests carried out in boreholes are

artificially induced hydraulic gradients as forcing for fluid

flow. Hydraulic tests provide data on the hydraulic proper-

ties and the nature of aquifers and the permeability struc-

ture of the upper crust including the depth variation of

permeability. Hydraulic well tests also have the potential to

identify and locate faults or fault systems in the crust. The

tests can quantify the hydraulic conductivity of the faults if

it is different from that of the fractured rock matrix. Fault

conductivity can be lower than, equal to, or higher than

that of the rock matrix. Examples of low-permeability

faults have been presented in Stober & Bucher (2005a)

and Stober et al. (1999). Important and frequently over-

looked driving forces for fluid flow are tidal forces that pre-

vent fluids in fractured systems from ever becoming

stagnant and chemically inactive. We also illuminate the

effect of tidal forces on flow and reaction of fluids in frac-

tured basement rocks.

PERMEABILITY, SIGNIFICANCE INFRACTURED BASEMENT ROCKS

The permeability j [m2] of a volume of fractured crystal-

line rocks forming the upper continental crust relates to

the structure and connectivity of the fluid-filled pore space.

The permeability controls, together with the viscosity and

density of the fluid, the ability of the volume of rock to

conduct a fluid phase. Permeability is, together with the

fluid properties, a decisive parameter controlling fluid flow

in brittle rocks. It can be retrieved from the transmissivity

T [m2 sec�1] calculated from hydraulic test data. From the

experimentally derived transmissivity T, the hydraulic con-

ductivity K [m sec�1] can be computed. Quantifying T

and K from pressure – time data obtained in well tests

requires geological expertise and ideas about the geological

and hydraulic structure of the subsurface. The parameter K

is not measured data; it is ultimately the result of a model-

ing process.

The Darcy flow law describing fluid flow through porous

media can be written as

q~¼ � K½ � 1gq

rP ð1aÞ

where q~ is the specific discharge vector per cross-sectional

area [m3 sec�1 m�2] with the components qx, qy, qz, [K]

[m sec�1] the tensor of hydraulic conductivity, q [kg m�3]

the density of the fluid, g [m sec�2] the acceleration due

to gravity, and ∇P [kg m�2 sec�2 or Pa m�1] the vector of

the pressure gradient.

The one-dimensional form of Darcy flow is

q ¼ �K1

gqDPD1

ð1bÞ

with ΔP [Pa] the pressure difference along the baseline Δl[m].

In hydraulic tests, Q [m3 sec�1] corresponds to the rate

of pumping fluid from or injecting fluid into a wellbore. It

is a parameter that can be varied during the test. Q is

related to q in Eq. 1b by Q/A = q with A [m2] the cross

section perpendicular to flow. The imposed input signal Q

results in a pressure response ΔP by the conductive system

that can be measured with instruments.

Data analysis of hydraulic test data requires a hydraulic

aquifer model. For instance, during the radial flow period

in a homogeneous isotropic aquifer, the hydraulic conduc-

tivity K is proportional to the pumping or injection rate

and the inverse of the drawdown and the thickness of the

tested formation in a simple and ideal case. This means,

that the derived hydraulic conductivity K depends on the

chosen concept for aquifer characterization. The hydraulic

conductivity K, and hence also the permeability j, is not aparameter that can be measured directly and unequivocally

(like length with a measuring tape or a folding ruler).

© 2014 John Wiley & Sons Ltd, Geofluids, 15, 161–178

162 I . STOBER & K. BUCHER

The permeability j [m2] can then be derived from

hydraulic conductivity K:

j ¼ Klgq

ð2Þ

with the viscosity of the fluid l [kg m�1 sec�1 or Pa sec]

and K, g, j and q from above.

To better understand the significance and meaning of

both the hydraulic conductivity K and the derived perme-

ability j, it can be helpful to look behind the routine of

hydraulic well testing. How are such hydraulic tests per-

formed, how is the wanted parameter derived from the

measured pressure data, and what kind of compulsory

assumptions need to be made during data processing? This

analysis may help evaluating the strength and significance

of the parameter ‘hydraulic conductivity,’ but also disclose

its weaknesses.

Because the subject of interest in this special volume of

‘geofluids’ is the permeability structure of brittle continen-

tal crust, hydraulic tests in deep boreholes are of prime sig-

nificance. At deep drilling sites, additional monitoring

wells are typically not available. The deep well is simulta-

neously used as testing and as monitoring well.

There is a generous repertoire of hydraulic tests tailored

and fitted to the specific application, environment, and

problem definition (and budget; Stober 1986; Krusemann

& de Ridder 1991; Horne 1996). Some continuously

record the pressure directly in the tested section (at depth);

in other tests, pressure is measured relatively close to the

surface. In many tests, the water level in the wellbore is

measured as a proxy for pressure rather than the pressure

itself. If no pressure data are available from the depth of

interest, convoluted corrections must be applied to near-

surface pressure or water level data. The flow rate Q can be

varied in hydraulic tests in many ways. Fluid can be

pumped or injected at a constant rate during the entire test

period; however, pumping rates can also be varied stepwise

according to a detailed protocol. The flow rate may also be

continuously adjusted to maintain a constant drawdown or

recovery of the water level in the wellbore. In the follow-

ing discussion, we presuppose that all tests are performed

according to the state of the art, that all data correction is

precise and correct, and that data evaluation and analysis is

based on a hydraulic model that correctly reflects the geo-

logical situation underground.

Test duration varies from a few minutes to several weeks.

The test produces information on the hydraulic properties

of the subsurface to a certain distance from the wellbore.

This distance may be named radius of influence. The radius

depends strongly on the test duration, the type of test

method, and the hydraulic properties of the rocks. More-

over, the radius of influence also depends on the technical

details of the special testing tools and the testing proce-

dure.

The hydraulic properties of the rocks in the vicinity of

the wellbore are commonly altered by the drilling process

itself or by drilling mud and added chemicals. Short-term

hydraulic tests will not be able to ‘see’ beyond this alter-

ation zone near the wellbore.

An example of the consequences of hydraulic tests with

small radii of influence is the 4444-m-deep wellbore Urach

3 (SW Germany). Urach 3 was drilled, as a research bore-

hole for deep geothermal energy utilization. From 1604 to

4444 m below surface, the bore is in granite and gneiss of

the Variscan basement. The research borehole provided an

unusually long series of water table measurements over a

period of 13 years. These Urach 3 data have been used to

decipher the hydraulic properties of the basement, the

chemical composition of the deep fluid, and the interaction

of this fluid with the rock matrix (Stober & Bucher 2004).

Urach 3 has also been used for a large number of diverse

hydraulic tests studying the influence of test methods and

test duration on derived hydraulic properties. For example,

long-term injection tests have been repeatedly performed,

along with slug tests of short duration. A comparison of

data from the two different types of tests (slug test and

log-term injection test) shows that the hydraulic conduc-

tivity around the wellbore is higher than of the pristine

formation by a factor of 7 (Stober 2011).

Flow rate and pressure variations during injection tests

can be delimited to the capacity of the formation. Derived

parameters will then reflect the intrinsic properties of the

formation. If tests are run outside these conditions, frac-

ture apertures may artificially increase, and elastic reactions

of the formation alter the conductivity of the system dur-

ing the test. The derived parameters reflect the artificial

properties created by the test itself. Injection tests in the

Urach 3 research well, for example, induced elastic reac-

tions by the gneissic basement at well head pressures above

170 bar, resulting in a significantly increased conductivity

(Stober 2011). In tectonic active areas, injection tests may

reduce effective stress and frictional strength of fractures.

The induced stress reduction by shearing can result in an

irreversible and permanent increase in the hydraulic con-

ductivity (Evans 2005). Consequently, injection tests per-

formed at an overpressure that is too high for the tested

bedrock alter the natural hydraulic properties of the tested

rock volume. Hydraulic parameters derived from such tests

do not reflect the properties of the natural system.

Hydraulic tests may test the entire open hole or seg-

ments of it, or the section with perforated casing. Often,

packers or other technical tools are used to isolate certain

sections of the wellbore for testing. The tested sections

typically vary from several hundred to only a few meters.

The prime result of processing and analyzing the data of

a hydraulic well test is the so-called transmissivity of the

tested section and not directly the hydraulic conductivity

or the permeability. The transmissivity T [m2 sec�1] is an


Hydraulic conductivity and fluid flow 163

integral quantity characterizing the hydraulic properties of

the entire tested section. If the tested section H [m] is

homogeneous and isotropic, the hydraulic conductivity can

be computed from the determined transmissivity:

K ¼ T =H ð3ÞThe transmissivity of tested heterogeneous and anisotropic

stratified formations can be written as an integral over all

contributions:

T ¼Z H

0

Kdh ð4Þ

The hydraulic conductivity K [m sec�1] can be calculated

from the experimentally derived transmissivity T from

Eqs 3 and 4.

Fluid flow in brittle crystalline basement rocks occurs

along fractures and faults, and along lithological contacts

and other discrete 2D structures. The water-conducting

features form an interconnected network with characteristic

hydraulic properties (Mazurek 2000; Mazurek et al. 2003;

Stober & Bucher 2007a). Water-conducting structures can

be regularly or unevenly distributed. Consequently, at

some drilling sites, the data can be interpreted as hydraulic

response of a homogeneous isotropic aquifer. At other dril-

ling sites, other concepts may appear more appropriate.

The choice of the concept depends on the amount of avail-

able geological information. If not much is known about

the geological structure, the derived parameter can be cor-

rect if the structure is simple or incorrect if the structure is

more complex than has been recognized.

Such considerations are not made during planning of

hydraulic tests, and the hydraulic conductivity is computed

from Eq. 3. This practice may lead to serious misjudg-

ments of the hydraulic conductivity and permeability of

fractured rocks. The following considerations may illustrate

the problem: A hydraulic test performed in a 600-m-thick

open hole of a deep wellbore resulted in a measured trans-

missivity T = 2 9 10�4 m2 sec�1. The calculated hydraulic

conductivity (Eq. 3) T/H = K = 3.3 9 10�7 m sec�1

appears feasible for fractured basement rocks. However,

the main water inflow to the open hole can be restricted to

a 10-m-thick very strongly fractured fault zone. The result-

ing hydraulic conductivity for the fault zone is

K = 2 9 10�5 m sec�1, 60 times higher than the first

value. Strongly channeled fluid flow is not uncommon in

fractured basement. Such zones can be identified and local-

ized by geophysical well logging. The example shows that

errors and misinterpretations can increase with increasing

length of the tested section. On the other hand, the exam-

ple also shows that the selection of an invariant test length

can be very difficult and the result questionable. The high

K-value may characterize the conductive fracture zone.

The lower K-value, however, may not characterize the

entire system of fractured rocks and highly conductive fault

zones.

Geophysical well logging methods such as flow meter

logs, conductivity logs, and temperature logs help to better

characterize heterogeneous hydraulic properties of tested

formations. Prominent inflow zones can be identified and

separately tested using packer systems. The methods may

produce a quantitative vertical permeability profile. How-

ever, such elaborate and costly test designs are rarely used

(typically in the context of development of repositories for

nuclear waste, for example, Nagra 1985, 1992). Normally,

one relies on T/H data to derive hydraulic conductivity K

and permeability j. It is thus evident that conductivity K

and permeability j in the end do not represent true and

absolute values. The derived permeability j of the conti-

nental crust, for instance, at a given point of interest,

strongly depends on the details of the test design of the

hydraulic test used for deriving the parameter. It may be

useful and representative, but the contrary may be true as

well if the basement is hydraulically complex and heteroge-

neous and the design of the test is unable to adequately

resolve the heterogeneity, or the tests are inappropriate to

derive a meaningful permeability at a point in the bedrock.

However, we do not propose to give up measuring proper-

ties by experiments, but rather to highlight possible pitfalls

of the methods and of uncritically using published perme-

ability data for modeling projects.

The hydraulic conductivity K characterizing the hydrau-

lic properties of the bedrock has two components. Firstly,

it describes the properties of the fracture network

expressed by the permeability of the rock volume, and sec-

ondly, it depends on the properties of the fluid itself. This

follows from Eq. 2, which may be rewritten to highlight

this aspect of K and j.

K ¼ jgql

ð5Þ

It can be seen from Eq. 5 that hydraulic conductivity

data derived from well tests depend on the properties of

the fluid residing in the fracture porosity. If one is inter-

ested in the permeability of the tested rocks at a point of

interest, the density q and the viscosity l of the fluid at

that point must be known (see Eqs 1a,b and 2 and defini-

tion of the parameters there).

The fluid parameters q and l depend on the tempera-

ture, the pressure, the mineralization, and the gas content

of the fluid. This means that hydrochemical analyses must

be available from sampling points at depth, in addition to

depth and temperature data. The factor l/gq in Eq. 2

converts hydraulic conductivity to permeability. It has the

value 1.0 9 10�7 m sec at 20°C for pure water, 1.7 9

10�8 m sec at 180°C for pure water, and 3.1 9 10�8 m

sec at 180°C for 100 g kg�1 NaCl solution (using l and qdata from H€afner et al. 1985; Sun et al. 2008; Wagner &



Kretschmar 2008). Particularly, the viscosity of the fluid in

the example above is markedly different in hot and highly

mineralized fluid from that of pure water. Pumping tests in

deep thermal wells can produce hydraulic conductivity val-

ues up to 8 times higher than would be obtained under

cool low-salinity conditions, independent of the permeabil-

ity structure of the formation itself, just as a result of fluid

properties. Derived permeability may be erroneously high

if fluid properties are unknown or not adequately consid-

ered. The opposite effect can be recognized in injection

tests, which normally use cool low mineralized surface

water as injection fluid.

The correct conversion of K to j data requires fluid data

from samples representative of the tested section. Collect-

ing and analyzing fluid samples from deep wells, particu-

larly if the fluid contains high amounts of dissolved gases,

is technically challenging and expensive, and therefore,

fluid data are often not available.

The permeability of rocks can also be measured on drill

cores in the laboratory. It is difficult or impossible to cor-

rectly represent fractures, faults, and larger cavities in core

samples. The laboratory-measured permeability typically

characterizes the property of the unfractured rock matrix.

In general, it is significantly lower than the permeability of

large volumes of fractured and faulted basement derived

from well tests.

The above discussion of the parameters hydraulic con-

ductivity and permeability relates to well tests in deep

boreholes that are the prime source of K and j data. The

technical details of such tests and their cross-links to the

geology at the drilling site have specific consequences for

the derived data. The user of these data must be aware of

these aspects, evaluate data carefully, and not adopt perme-

ability data uncritically.

Long-term pumping tests in deep boreholes may be well

suited to detect the hydraulic properties of fractured base-

ment. Under favorable circumstances, long-term tests can

identify, locate, and characterize large-scale hydraulic struc-

tures (so-called hydraulic boundaries). An excellent exam-

ple is the highly developed pumping test in the 4-km-deep

pilot hole of the continental deep drilling project (KTB) in

Germany (Stober & Bucher 2005a). The test produced a

reliable value for the hydraulic conductivity of homoge-

neous fractured continental basement at 4 km depth

K = 4.07 9 10�8 m sec�1, which results in a characteristic

permeability j = 1.18 9 10�15 m2 given the fluid proper-

ties at depth (120°C, TDS = 62 g kg�1, l = 0.27 9 10�3

Pa sec, q = 952.8 kg m�3; mol weight NaCl = 58.44

g mol�1). The fluid temperature and the salinity have been

taken from Stober & Bucher (2005a), the dynamic viscos-

ity of the fluid at 120°C and 1 molal salinity comes from

Kestin et al. (1981), and the density of a 1 molal salt solu-

tion at 120°C comes from Rogers & Pitzer (1982). This

value of the permeability j = 1.176 9 10�15 m2 may be

considered a representative value for fractured continental

crust at about one-third of the distance from the surface to

the brittle ductile transition zone in moderately warm crust

(MOHO temperature � 650–700°C). The value is in

excellent agreement with permeability data from continen-

tal crust published by Ingebritsen & Manning (2010).

Their permeability data have been compiled from conti-

nental crust worldwide and have been derived by various

geophysical tools and methods but not from hydraulic well

tests. Our j = 1.18 9 10�15 m2 value is closer to the new

j versus depth fit of Ingebritsen & Manning (2010) than

to their previous fit (Ingebritsen & Manning 1999).

Hydraulic testing of the KTB wellbore discovered a

hydraulic boundary with a lower conductivity than the

fractured basement. The seal appears to be located at a dis-

tance of several hundred meters from the wellbore at 4 km

depth. The hydraulic boundary was identified as a fault and

correlated with a major fault zone known from surface

geology. Similar ‘water-tight’ hydraulically prominent faults

have been identified at other localities in the central Euro-

pean crystalline basement (Stober et al. 1999). Both exam-

ples show that prominent fault systems in the basement are

not necessarily characterized by a higher conductivity than

the surrounding fractured basement.

PERMEABILITY AND FLUID FLOW IN THECRUST

Hydraulic tests in wellbores to 5000 m depth worldwide

reveal a remarkable variation of hydraulic conductivity of

the crystalline basement from 10�14 to 10�4 m sec�1. The

upper 1000 m are generally characterized by the higher

values but also by a greater variance (Stober 1996; Stober

& Bucher 2007a). The mean variation of the hydraulic

conductivity decreases rapidly with increasing depth. The

decrease in the permeability j with depth has been derived

from well tests in the basement of SW Germany, NE

France, and N Switzerland (Stober & Bucher 2007a,b)

and can be described by

log j ¼ �1:38 log z � 15:4 ð6Þwith z the depth in km and j in m2. The described

decrease in the permeability with depth (Eq. 6) is mainly

controlled by the properties of biotite gneiss and other

metamorphic rocks because test data for greater depths

were mostly available from wellbores in gneiss. We suspect

that the decrease in the hydraulic conductivity in granitic

basement would be slower with depth, because of contrast-

ing mechanical properties of mica and feldspar (see above

and in Stober & Bucher 2007a). At our studied sites, gra-

nitic basement has a higher permeability at a given depth

and tectonic environment than gneissic basement (given

that granite is dominated by feldspar and quartz and gneiss

typically contains significantly more mica than granite).



Ingebritsen & Manning (1999) in their geophysical study

of terrestrial heat flow derived a surprisingly similar power-

law function for the permeability decrease with depth.

Their function predicts at 4 km depth in crystalline base-

ment that log j = �15.9 m2. Ingebritsen & Manning

(2010) compared two k–z relations: (i) for ‘tectonically

active’ crust, that is, their 1999 curve; and (ii) another

curve for ‘disturbed crust’ that included hypocenter migra-

tion, EGS, and other sources of data. The new function

predicts a permeability for ‘disturbed crust’ at 4 km depth

(KTB pilot hole) of log j = �13.5 m2 (Eq. 12 Ingebritsen

& Manning 2010) in contrast to our derived value of log

j = �14.9 m2 at the KTB site.

The decrease in the hydraulic conductivity and perme-

ability with depth has, to our knowledge, never been

reported and demonstrated for a single discrete deep

borehole (>1000 m) in the crystalline basement. In deep

wellbores, adequate hydraulic tests at different depths over

sufficiently long test intervals consistent with the require-

ments of a correctly chosen REV are not normally avail-

able. Urach 3 is an exceptionally deep research bore in

fractured crystalline basement of SW Germany. Several

hydraulic tests at different depths and a total test interval

of more than 2000 m produced unique hydraulic data.

The basement belongs to the so-called Moldanubic zone

of the central European continental crust. The basement

consists of mafic biotite–hornblende gneiss (amphibolite),

migmatitic gneiss, quartz–diorite, biotite–cordierite gneiss,

and other high-grade metamorphic rocks. The unaltered

high-grade rocks contain zones of hydrothermal alteration.

The permeability of the basement is related to a fracture

network of variable orientation. There are open fractures

functioning as water-conducting structures and sealed

impervious fractures. The fracture density is about 2–3

open fractures per meter (Dietrich 1982). In the years

1978–1979, the wellbore was drilled to 3334 m (bottom

hole at that time) with a 14-m-thick open hole section.

The casing has been perforated at different sections, and

the accessible rock has been hydraulically tested. The well-

bore has been deepened from time to time (for different

reasons) in the following years. At each increment, the

newly accessible rocks have been hydraulically tested. In

1992, the wellbore reached its present depth of 4444 m

(Stober 2011). The incremental drilling history of the

Urach 3 bore offers a unique opportunity to evaluate the

variation of the hydraulic conductivity with depth

(Fig. 1).

Transmissivity data have been derived from pumping and

injection tests at 1.77, 3.30, 3.35, 3.88, and 3.97 km depth,

respectively. The hydraulic conductivity (T/H) of fractured

basement gneiss decreases systematically with depth z. At

3.97 km, the hydraulic conductivity is about 2000

times lower than at 1.77 km. The depth z – log (T/H) data

follow a linear trend (Fig. 1). The log (T/H) versus log z

data can be regressed to a linear equation:

logðT =H Þ ¼ �8:748 z � 4:465 ðR2 ¼ 0:976Þ ð7ÞThe log-linear Eq. 7 can be converted to Eq. 8 express-

ing the depth dependence of the permeability in crystalline

basement rocks (here using data for a fixed salinity of 1

molal NaCl, 120°C and 300 bar).

log j ¼ �8:748 z � 12:0 ðR2 ¼ 0:976Þ ð8Þlog j ranges from �14.24 at 1.77 km depth to �17.56 at

4 km depth. Note that the permeability of gneiss basement

at the Urach 3 site is about 400 times lower than that of

the basement at the KTB site at the same depth of 4 km.

In general, the hydraulic conductivity of granite is higher

than that of gneissic basement in areas of strong and

young deformation (e.g., the three-country corner of

Germany–Switzerland–France). In tectonically inactive

areas, fracture density and, consequently, the hydraulic

conductivity of large volumes of granite can be very low

(Stober & Bucher 2007a).

With increasing depth, the vertical stress component

usually exceeds the horizontal stress components (Brown

& Hoek 1978) with the result that the open water con-

ducting fracture system changes from predominantly hori-

zontal to mostly vertical orientation. Vertical fluid flow and

fluid exchange becomes dominant at depth, particularly in

areas with significant topography (Fig. 2). The steep fluid-

conducting fractures support deep fluid circulation systems

Log (T/H) [m sec—1]

Dep

th [k

m]

Fig. 1. Log (T/H) data from well tests at five different depths in the Urach

3 borehole in gneiss of the Black Forest crystalline basement.



(Stober et al. 1999; Stober & Bucher 2004; Bucher et al.

2012). Hydrothermal alteration zones along vertical frac-

tures document the existence of significant vertical perme-

ability (Lee et al. 2011; see also section on reactive fluid

flow below).

One notable exception has been described from the

Blancket well (Australia), where the vertical pressure com-

ponent becomes lower than the horizontal component at

depths below 1800 m. Consequently, fractures and faults

have a preferred horizontal orientation in the 4000-m-deep

stimulated geothermal Habanero 4 well (Bendall et al.

2014).

Hydrochemical and isotope data suggest that thermal

spring waters represent upwelling deep waters (Chebotarev

1955; T�oth 1962; Stober 1995; Allen et al. 2006; Bucher

et al. 2009). The conclusions from chemical data are sup-

ported by the evidence from numerical modeling (Forster

& Smith 1989; Kukkonen 1995). Topography-driven hot

water flow from several thousand-meter-deep sources has

been reported (T�oth 1978; Bucher et al. 2009).

The circulation systems depend on the existence of steep

fault structures that must merge into high-conductivity

fault zones allowing for rapid upward flow of deep water

(e.g., Baden-Baden; Stober 1995; Sanner 2000; Fig. 2).

The final temperature of the deep fluid reaching the sur-

face depends, among other factors, on the flow velocity,

which is proportional to the hydraulic conductivity for a

given hydraulic gradient. Typically, major fault zones with

vertical displacements of hundreds or thousands of meters

tend to have an inner, low-permeability central zone man-

tled by high-permeability rims on both sides (Mazurek

et al. 2003).

Topography-driven deep water circulation in the Black

Forest region (Germany) results in the outflow of warm

saline deep water into the freshwater aquifer of the Quater-

nary fill of the drainage valley (Kinzig valley). The natural

contamination of the freshwater by deep salty fluids can be

traced as a NaCl plume (Ohlsbach Plume) for a long dis-

tance into the Quaternary gravel aquifers of the river Rhine

plain (Stober et al. 1999). The deep saline warm water

flows upward along a steep fault system forming the

boundary of the Black Forest topographic high and the rift

system of the river Rhine valley (Fig. 3). Although the rift

escarpment represents a generally extensional environment,

strike-slip movements along the faults form fine-grained

fault gauge, which clogs permeability, thus forming

hydraulic barriers. However, many of the Rhine rift bound-

ary faults are highly transmissive and could channel fluid

flow (e.g., Bruchsal and Landau; Aquilina et al. 2000;

B€achler 2003).

In the case of the Ohlsbach plume (Fig. 3), deep water

is forced to flow upward by the hydraulic boundary fault

through fractured granitic rocks that have similar conduc-

tive properties as the area of recharge. The flow velocity is

slow so that the water reaches the surface with only 27°Calthough the reservoir at >3000 m depth has a tempera-

ture of close to 100°C (Stober et al. 1999). If upwelling

deep water from similar reservoirs follows a highly perme-

able fault zone such as in Baden-Baden, the outflow tem-

perature is as high as 68°C although the reservoir

temperature may be as high as 150°C. The water is charac-

terized by a discharge temperature that reflects the perme-

ability structure of the faults and the basement as

described by theoretical models of fluid flow in fault zones

(L�opez & Smith 1995).

Hydrothermal circulation systems, ascent and descent

channels alike, require open highly permeable steep frac-

ture systems. Ascent channels must be extremely conduct-

ing structures permitting high flow rates in order for hot

water to reach the surface environment at high tempera-

Fig. 2. Schematic water flow systems in a crystal-

line basement area with moderate topography

(e.g., Black Forest basement). The permeability

contrast between granite and gneiss has a

strong control on the flow paths, in addition to

topography and fault structures. The figure also

illustrates changing water composition along

flow.



ture and appear as a hot spring. The ascent flow may fol-

low zones of strongly fractured granite (Fig. 2) or low-per-

meability fault zone (Fig. 3). Driving force for the

circulation is topography in both examples. In areas lacking

topography, or in the absence of suitable steep high-

permeability structures, deep-reaching vertical circulation

systems may not develop, and natural hot springs will not

be present.

REACTIVE FLUID FLOW IN THE CRUST ANDITS EFFECT ON PERMEABILITY

At depth below the deepest borehole that has been

hydraulically tested, the permeability of the continental

crystalline crust can be derived from other data and obser-

vations including geothermal data (Ingebritsen & Manning

2010).

An important source of information for understanding

permeability and permeability evolution with time in dee-

per parts of the brittle crust is the preserved effects of fossil

fluid–rock interaction in surface outcrops of fractured rocks

from the zone between the deepest wells to the brittle

ductile transition zone. The temperature range where these

structures formed is about 200–400°C. Fluid flow along

fractures in crystalline basement rocks is generally accom-

panied by chemical reactions between the aqueous fluid

and the rock exposed along the fracture. The reactions

occur because the fluid is rarely in chemical equilibrium

with the mineral assemblages of granites and gneisses pre-

dominantly present in the continental crust (Garrels &

Howard 1959). At the generally relatively low temperature

above the ductile–brittle boundary (<400°C), reactions

involving primary silicates are generally slow (Browne et al.

1989; Brantley 2004; Yadav & Chakrapani 2006) and

most high-grade minerals are not stable in the presence of

H2O at the prevailing conditions. Fluid–rock reaction (ret-

rograde metamorphism) occurs from the ductile–brittle

boundary to the surface in fractured basement that ranges

in temperature from about 400°C down to ambient tem-

perature. Hydration of primary high-grade assemblages is

the major chemical process. Newly formed clay, zeolite,

Fe-hydroxide, and other alteration products are common

on fractures studied at drill cores from deep boreholes and

from tunnels (Stenger 1982; Bauer 1987; Borchardt et al.

1990; Borchardt & Emmermann 1993; Lindberg &

Siitari-Kauppi 1998; Iwatsuki & Yoshida 1999; Weisenber-

ger & Bucher 2010). This shows that fluid–rock interaction

and associated mineral dissolution and precipitation along

the fractures contribute to the variation of the permeability

with time (Moore et al. 1983; Polak et al. 2003).

Mineral coats along the fractures develop from the

instant of fracture formation and its saturation with aque-

ous fluid until the fracture becomes sealed by reaction

products. It is difficult to relate veins of observed second-

ary minerals in drill cores to the flow-active fracture poros-

ity and to the quantitative permeability history of the

crust. However, these structures relate to fossil flow and

reaction systems and the observations can be regarded as

analogs for present-day processes progressing in the deeper

parts of the upper crust.

There is abundant evidence for permeability-relevant

reactions in deep fractured rocks. The effects of the reac-

tions can be easily recognized in so-called reaction veins

(Fig. 4). The dissolution and precipitation processes may

proceed at different time-varying rates. Reaction veins have

a structural component caused by rock deformation and a

chemical reaction component that follows from irreversible

reaction of the advective fluid with the exposed rock. Both

components have consequences for the permeability of the

fractured system and its variation with time.

Simple monomineralic veins may form by a crack–seal

mechanism (Durney & Ramsay 1973; Ramsay 1980; Ankit

et al. 2013). These veins have a predominantly deforma-

tional component and include quartz veins in quartzite (or

granite) and calcite veins in limestone and marble. Reactive

flow may dissolve components from the exposed rocks,

Fig. 3. Section from the Rhine River rift valley to

the Black Forest basement along the Kinzig

River valley. Horizontal scale of section about

20 km. Recharge water from the Black Forest is

channeled along a fault system to the surface

and does not reach the Rhine graben structure.

For more details, see Stober et al. (1999).



import dissolved components from external sources, and

precipitate new minerals in the fractures that were not ini-

tially present. These veins result from a fracture–reaction–

seal mechanism (Bucher-Nurminen 1989). Examples

include serpentine veins in peridotite (Fig. 4), talc veins in

peridotite (Fig. 5), chlorite veins in biotite gneiss and in

garnet amphibolite (Fig. 6), and tremolite veins in marble

(to name a few examples). The structures shown on

Figs 4–6 contain a wealth of qualitative information on the

development and time dependence of permeability in the

brittle crust. Some of the field-deduced conclusions are

summarized in the figure captions.

These examples suggest that vein growth starts with an

initial brittle fracture, followed by a reaction period during

which permeability increases, and then a period of decreas-

ing permeability until the vein becomes impervious and

fluid flow stops. Thus, permeability follows a time evolu-

tion similar to the porosity wave proposed for the lower

crust (Connolly & Podladchikov 2007). The schematic

permeability–time relationship shown in Fig. 7 character-

izes a single vein or a single fluid-conducting structure

(e.g., shown in Fig. 6). The permeability evolution path

(Fig. 7) shows the transient local effect of a single fracture

on top of the original background permeability. After seal-

ing the fracture, the system returns to the background per-

meability, which represents the sum of all effects of all

flow-active fractures currently contributing to the perme-

ability of the considered volume of rock. If many veins are

active simultaneously, transient permeability increases and

then, with reaction progress, permeability decreases. How-

ever, if vein formation is not strictly synchronous in an

area, the single fractures contribute statistically to the back-

ground permeability.

Outcrops on Vannøya (northern Norway) representa-

tively show the sequence of a crack–reaction–seal cycle

(Fig. 8). First, fluid-conducting fractures were formed in

the green mafic igneous rock. The distinctive green–brown

reaction front resulted from reaction of a CO2-rich fluid

with the green rock. These reactions destroyed all Fe-bear-

ing minerals and carried away dissolved iron, leaving a

brown albite–carbonate rock. It is not a redox reaction

front (e.g., Yamamoto et al. 2013) but rather a carbon-

ation front (Priyatkina et al. 2011).

Hydrothermal alteration reactions may, in their simplest

form, be plain hydration reactions. No components dis-

solved in the fluid are lost or gained in such reactions.

Hydration reactions consume H2O and cease if water sup-

ply to the reactive fracture ends. Consumption of H2O

passively increases the total mineralization of the fluid

(TDS). Because of hydrothermal alteration (retrograde

metamorphism) of upper crustal rocks and the associated

H2O consumption, deep fluids are generally highly saline

(Frape & Fritz 1987; Edmunds & Savage 1991; Stober &

Bucher 2005b; Bucher & Stober 2010). The salinity is typ-

ically much higher than that of seawater (up to NaCl satu-

ration). Hydration reactions also desiccate the ductile

Fig. 4. Serpentine reaction veins in peridotite (Erro Tobbio mantle, Italy).

The serpentinization of peridotite occurred along symmetrical reaction veins

along brittle fractures. During the active period of serpentinization, the per-

meability of the fractured rock permitted flow and reaction of aqueous

fluid. Flow and reaction ceased because of self-sealing of the fractures, pre-

venting the peridotite (brown crust) from being completely serpentinized.

Later, the fractures became completely sealed, and the vein system does

not contribute to the near-surface permeability of the partly serpentinized

peridotite.

Fig. 5. Talc veins in peridotite (Vlisarvatnet, Norway). Low-temperature

hydration of peridotite (coarse-grained spinel harzburgite with brownish

weathering rind) produced white silvery zones of talc along open brittle

fractures that represent fluid-conducting structures. Note the vertical orien-

tation of the fractures and the large aperture of these young structures.

High content of dissolved silica in the hydrothermal fluid resulted in talc

formation rather than serpentinization (Fig. 4). The open fracture system

contributes to the present-day permeability of the rocks (and could be

measured with well tests).



lower continental crust (Frost & Bucher 1994). Evidence

for active desiccation of the lower crust is the presence of

high-temperature halite in eclogites (Markl & Bucher

1998).

Plagioclase is the most abundant mineral of the conti-

nental crust. This feldspar mineral is a mixture of a Ca

component (anorthite) and a Na component (albite). At

low-temperature hydrothermal conditions, Ca-bearing ver-

sions of the mineral cannot be in equilibrium with H2O,

and it hydrates to clay or various zeolite minerals such as

stilbite, laumontite, and others depending on the tempera-

ture (Nishimoto & Yoshida 2010; Weisenberger & Bucher

2010) and pressure (depth). Low-temperature hydration of

the anorthite component of plagioclase can be written as

Anorthiteþ 5 quartzþ 7H2O

¼ stilbiteðCaAl2Si7O18 � 7H2OÞ ð9ÞThe mineral name stilbite is used here for a Na-free end

member component stilbite–Ca. It forms at very low tem-

perature (<120°C). Quartz on the reactant side of Eq. 9

can be taken from the minerals exposed on the fracture

surface or from dissolved SiO2aq in the fluid. In the pro-

(A)

(B)

Fig. 6. Reaction veins in Caledonian amphibolite-grade gneiss: (A) Early

chlorite vein in biotite gneiss (Hammerfest, Norway). Biotite gneiss has

been chloritized along a brittle fracture that served as fluid conduit for

hydrothermal fluid. The open fracture has been sealed later by epidote and

became inactive. Young brittle fractures are open water-conducting struc-

tures and represent the present-day permeability of the rock. The late

structures (blue arrows) control the hydraulic conductivity detected and

measured by well tests. (B) Chlorite vein in garnet biotite gneiss (Torsnes,

Kvaløya, Norway). Chloritization occurred along fractures by interaction of

advecting hydrothermal fluid with the primary gneiss assemblage. Most of

the fractures are completely sealed by solid reaction products. One of the

fractures is only partly sealed. The visible fracture porosity contributes to

the present-day permeability of the rocks (detectable by well tests).

Fig. 7. Schematic graph showing the time-dependent permeability devel-

opment related to a single fracture that serves as a fluid conduit for reactive

fluids. This simple pattern can be modified in many ways, for example, by

multiple fracturing, time-dependent aperture variations due to complex

extension and shearing, and by externally controlled changes of fluid com-

position. In addition, the permeability of the large-scale permeability is con-

trolled by the combined result of j–t curves of a large number of fractures

of different orientations.

Fig. 8. Vein system in mafic rocks (dolerite on Vannøya, Norway). The

structures suggest that first a fracture system opened, then advecting fluid

reacted with the mafic igneous rock (green) producing an albite–calcite rock

(light brown) before the fluid conduit was finally sealed by brown (and

white) carbonate in the central part of the veins. The structures support the

proposed j–t curve shown on Fig. 7.



cess, the albite component of the plagioclase recrystallizes

as pure albite. Laumontite, another Ca-zeolite, forms from

hydration of plagioclase at higher temperature (~150–250°C) corresponding to a depth range of 6–10 km. Its

formation reduces hydraulic conductivity and may effi-

ciently desiccate the fracture system (Fig. 9).

Many primary igneous minerals hydrate to zeolites and

other hydrous minerals when the igneous high-temperature

rocks pass through the brittle upper crust during exhuma-

tion before reaching the erosion surface. Their production

on fractures typically involves a large volume increase rela-

tive to the original igneous rock. The process efficiently

seals the water-conducting fracture, and the volume

increase of the solids can accommodate significant fracture

extension. The large volume increase of the solids in

hydration reaction seals the fracture porosity and stops

veins from growing. The lifetime of high-permeability con-

ditions permitting fluid flow is difficult to quantify. How-

ever, the permeability of sealed veins may not be higher

than that of the rock matrix.

Generally, zeolites are widespread and common in frac-

tured continental basement. They have been reported to

occur on fractures and fissures encountered during road

and rail tunnel construction in the Alps (Armbruster et al.

1996; Weisenberger & Bucher 2010), in deep drillholes in

the continental basement (Stober & Bucher 2005b) and in

drilled young granitic plutons (Yoshida et al. 2013). An

example of zeolite-bearing veins is shown in Fig. 9. Pink

Ca–zeolite laumontite has formed along vertical fractures

through massive coarse-grained gabbro of the Seiland igne-

ous province in northern Norway. The exposed fractures

are efficiently sealed by zeolite that formed according to a

reaction analogous to Eq. 9 above. The presence of lau-

montite instead of stilbite as fracture filling (and other pet-

rologic details) suggests that the hydration reaction

occurred at a temperature of about 250°C. Thus the

hydration veins in fractured basement displayed on Fig. 9

represents a situation at about 10 km depth, assuming a

geothermal gradient at the time of hydrothermal vein for-

mation of about 25 K km�1. Note also that the orienta-

tion of the fractures is nearly vertical as a result of

predominance of vertical stresses over horizontal stresses at

this depth (this is also the case shown in Fig. 5). The

hydration reactions progress in two steps: A first step dis-

solves the reactive high-temperature minerals such as igne-

ous plagioclase, pyroxene, and olivine. The dissolution step

produces porosity and increases related permeability.

Plagioclase conversion to albite and zeolite has produced

initially up to 15 vol.% porosity in fractured granites of the

Alps (Weisenberger & Bucher 2010). In a second step, the

created secondary porosity is replaced by solid reaction

products that eventually seal the open fractures. Fluid–rock

reactions in fractured granites thus first enhance permeabil-

ity that originally has been created by mechanical fractur-

ing of the granite. Increasing amounts of solid reaction

products gradually decrease permeability of the porous,

fractured wall rock rocks (Fig. 7).

Hydration reactions consume H2O from the fluid and

create a local pressure depression inducing of fluid flow

toward the reaction site. In the Urach 3 borehole, vertical

downward migration of fluid has been associated with

progressing hydration reactions at depth (Stober & Bu-

cher 2004). H2O consumption at depth may not be fully

compensated by downward fluid flow because of low per-

meability. Consequently, a measurable vertical hydraulic

potential gradient in the Urach 3 bore, with decreasing

potential with increasing depth, is enhanced and sup-

ported by low and decreasing hydraulic conductivity with

depth. Thus, downward fluid flow can be small despite

favorable steep fracture orientation at depth simply due to

low conductivity of the fractures. Deep fluid circulation

systems may not develop in such low-permeability envi-

ronments.

There are few direct observations of changing permeabil-

ity of crustal rocks with time. This is a consequence of the

timescale of the fracture–reaction–seal processes in the

crust. The effects cannot normally be detected in well tests.

Even if well tests could be repeatedly performed in 5-km-

deep boreholes over periods of 20–30 years, well detoria-

tion and other technical effects would probably obliterate

the signals from the undisturbed ground.

Fig. 9. Laumontite (Ca–zeolite) clogs vertical fluid-conducting fractures in

gabbro (Langfjorden, Norway). The zeolite forms from hydrothermal alter-

ation of primary labradorite (plagioclase) of the gabbro. The large volume

increase in the solids in the reaction efficiently seals the vertical fracture sys-

tem. The vein formation occurred at about 250°C and 10 km depth during

the late Caledonian formation of the Langfjorden–Repparfjorden fault

system.



In tectonically active areas, water level measurements

and other data collected after earthquakes have been used

to show that active fracturing temporarily increased perme-

ability. The healing of the created deformation structures

occurred relatively rapidly, within the timescale of observa-

tion. Generally, the observed and documented permeability

changes refer to mixed crust (basement and sediments) or

ocean island crust, and not the inactive normal continental

crust (Elkhoury et al. 2006; Kitagawa et al. 2007; Cappa

et al. 2009; Xue & Cast 2013).

Historical reports on the behavior of temperature and

discharge of thermal waters in geothermal or mineral water

spas indicate variations with time. One example is the spa

Badenweiler in SW Germany. This large and luxurious spa

was built and used by the Romans about 1800 years ago.

The buildings were once heated by hot water using a

sophisticated thermal heating system. Today, discharge and

temperature of the hot springs are insufficiently low for spa

operation. This indicates that the permeability structure

has changed on the timescale of hundreds of years (Filgris

2001).

FLUID FLOW AND PERMEABILITYSTRUCTURE OF THE UPPER CRUST

The Darcy flow law (Eq. 1b) implies that fluid flow ceases

if the pressure gradient disappears. The major forces driv-

ing fluid flow in the crust are thermal disequilibrium and

topographic relief. If the driving force ΔP (Eq. 1b)

approaches zero, fluid flow stops. The fluid becomes stag-

nant. The stagnant fluid resides in the fracture pore space

and chemically interacts with the rock matrix at a very slow

rate. The fluid interacts with solids and pores by diffusion.

In low-permeability near-surface rocks such as clay and

shale, with permeability lower than 10�20 m2 (10 nD), flu-

ids are considered stagnant, and the Darcy flow law cannot

describe fluid transport because at very low permeability

and feasible pressure gradients, the flow–force relationship

is not linear (Bear 1979).

In fractured crystalline basement rocks with a typical per-

meability of 10�15 m2, fluid flow stops if the driving pres-

sure gradient vanishes, for instance, because the

topographical relief has been eroded and removed. There-

fore, the fluid in the fracture pore space of basement of flat

vast plains has been considered essentially stagnant.

This view ignores a very efficient driving force for fluid

flow, the Earth tides. Two times every day a tidal wave

moves through the Earth crust. The rise and fall of the

Earth surface amounts to some tens of cm (Emter et al.

1999). This is because the Earth is not a rigid body; it

reacts elastically to the gravitational forces of moon and

sun. The moving layers of rocks exert compressional and

extensional forces on each other. The pore space of the

rocks also deforms elastically in reaction to tidal forces,

with the result that the fluid residing in the pore space

experiences an everlasting alternation of compression and

extension. The consequence of the ever-changing tidal

forces can be observed as water table fluctuations in bore-

holes. The magnitude of the fluctuations depends on the

position of the sun and the moon (e.g., Bredehoeft 1967;

Evans & Wyatt 1984; Hsieh et al. 1988). Therefore, the

highest amplitudes occur at full moon and new moon. In

the fractured crystalline basement at the deep drilling site

Urach 3, water table fluctuations related to tidal forces of

up to 20 cm per day have been measured (Stober 2011;

Fig. 10).

It can be concluded that the observed fluctuations of the

water table in the wellbore indicate the reaction of a very

large volume of deep water. This conclusion is in accord

with the very low fracture porosity of about / = 0.002 and

the very low compressibility of deep fluid cw = 5.3 9

10�10 Pa�1. A pressure change of ΔP = 1 bar results in a

relative change of the volume of water of DVw/Vw = 5.3 9

10�10 and relative to the crust of DVw/Vrock = 1.1 9

10�12. This implies that the fracture pore space is intercon-

nected on a large scale and the crust reacts in a hydrauli-

cally coherent fashion.

Long-term hydraulic tests in deep boreholes confirm

that a large volume of fluid hydraulically reacts and that

the fracture porosity of the basement is an interconnected

network of water-conducting structures. The 1-year-long

pumping test at the 4000-m-deep research drillhole (KTB-

VB: see above) in SW Germany extracted a total of

22 300 m3 saline thermal fluid (Erzinger & Stober 2005)

from the open hole at 3850–4000 m at a rate of

0.5 l sec�1 and later 1.0 l sec�1 (Fig. 11). The composi-

tion of the thermal fluid remained constant during this

time (Stober & Bucher 2005a). Given the porosity of

0.5%, the homogeneous volume of fluid originates from a

cylinder of fractured rocks with a radius of about 100 m

around the wellbore. The hydraulic signal reaches far

beyond this cylindrical volume around the bore. The

extracted fluid must be replenished from a much larger

volume around the borehole.

In regions with high topographic relief, deep flow sys-

tems develop in the crystalline basement. Temperature

profiles in very deep boreholes can be used to estimate

the water circulation depths. For instance, in the 5000-m-

deep geothermal well, GPK-2 in Soultz-sous-Forets near

Strasbourg, France, (Genter et al. 2010) efficient water

circulation reaches to 3700 m depth, which is 2300 m

into the crystalline basement. Below 3700 m, the geother-

mal gradient is identical to the regional average

(~27 K km�1). Above this depth, water temperatures indi-

cate advective heat transfer by upwelling hot deep water.

Finally, in the near surface, fluid temperature decreases

very rapidly, resulting in a very high geothermal gradient

(>100 K km�1). Thermal signatures of advective fluid



flow may be absent in some areas because permeability of

the basement is low or because pressure gradients gradu-

ally diminish (with the exception of the cyclic tidal

forces).

Temperature profiles of km-deep boreholes, data, and

observations from hydraulic tests and tidal water level fluc-

tuations all consistently show that fluids in the continental

crust occupy an interconnected communicating pore space

(A)

(B)

Fig. 11. Well test design and test response data from hydraulic tests in fractured continental basement in the KTB pilot hole at 4 km depth (Oberpfalz, Ger-

many).

Fig. 10. Water table fluctuations in the Urach 3

borehole caused by tidal forces.



that permits fluid flow on a large scale, provided that an

appropriate driving force exists in addition to the tides. In

the examples above, topography drives the fluid flow at So-

ultz-sous-Forets and pumping, an anthropogenic force,

drove fluid flow at KTB. Hot springs discharging highly

mineralized waters from crystalline basement for 2000 and

more years are clear evidence for deep fluid circulation and

an interconnected fracture network providing sufficient

permeability for fluid flow in the brittle crust.

Fluids in the crust are daily and perennially in flow

motion driven and pumped by tidal forces. This motion

mixes deep fluids, and chemical reactions can be reacti-

vated. The small-scale movement of aqueous solutions in

and out of microcracks, and the associated mixing pro-

cesses have a significant effect on the kinetics of hydrother-

mal reactions. Deep waters in the brittle upper crust are

never truly stagnant because tides keep them permanently

in movement. The main effect of fluid movement caused

by tidal forces is not its contribution to advection or bulk

flow, but on reaction kinetics, and local fluid composition,

and thus on the timing of the permeability increase and

decline.

The Earth tidal forces press the fluids into fine hairline

cracks and dead-end pores and recover it from these pores

during the decompression cycle. This cycling has significant

geochemical consequences. The flow mixing is much more

efficient than diffusional mixing. It maintains hydrogeo-

chemical reactions going at significant rates. The tidal flow

generates chemical reaction fronts along water conducting

fractures.

Extension of Darcy flow concepts to the lower crust is

problematic. The lower crust is normally dry, with any free

H2O/CO2 fluids removed by retrograde reactions (e.g.,

Frost & Bucher 1994). In an unusually active crust, pro-

grade dehydration of H2O-bearing minerals produces a

free aqueous fluid, and these fluids must be removed from

the lower crust. The exact transport mechanism is

unknown. Nonetheless, Connolly and Podladchikov

(2015) and Weis (2015) invoke a linear force–flux equa-

tion to quantitatively describe fluid removal from ductile

lower crustal rocks.

SUMMARY AND CONCLUSIONS

Permeability j and hydraulic conductivity K of the brittle

and fractured upper continental crust can be derived from

pressure–time data obtained from hydraulic tests in deep

boreholes. A handful of bores worldwide have produced

permeability data for the crystalline basement to depths of

about 4–5 km below surface. Below this depth accessible

to direct measurement of transmissivity, j and K data can

only be derived from indirect geophysical observations

(e.g., seismicity propagation rates). However, it appears

that permeability gradually decreases to the depth of the

brittle–ductile transition zone at about 15 km depth and

350°C in normal stable continental crust. As brittle water-

conducting structures become fewer toward the transition

zone, the representative elementary volume (REV, Bear

1979) increases toward the ductile lower crust, where

long-lived water-conducting structures do not exist. In the

ductile lower crust, free fluid exists only temporarily (Frost

& Bucher 1994; Yardley & Bodnar 2014).

For the uppermost 4–5 km of the brittle continental

crust, where direct transmissivity (T) data can be measured

in deep wells, the conversion of the T to j and K (=T/H)

data is not simple. The conversion to permeability requires,

for instance, the composition and properties of the fluid

occupying the fracture pore space. These data require rep-

resentative water samples. The sampling requires sophisti-

cated and expensive sampling techniques. Derived T data

from well tests are strongly influenced by the type of test,

its duration (long, short), and the technical details of the

test design and structure around the borehole. Injection

tests may alter the undisturbed permeability structure of

the formation, reversibly or irreversibly. A critical parame-

ter for converting T to K (j) is the test length of the

hydraulic test. It is often thought or assumed that perme-

ability characterizes the rock formation at a certain spot in

the crust. However, it may vary strongly with the test

length (H). A meaningful permeability refers to a represen-

tative test length, correctly representing the entire system

of rocks and fractures with all their properties including

fracture density, aperture, connectivity, the number, distri-

bution and discharge of water inflow points, and other

properties and structures relevant to fluid flow. The repre-

sentative test length may be very large in little deformed

basement. Thus, the characteristic permeability refers not

to a spot in the basement but rather to a potentially large

volume (e.g., several hundred meters). The representative

test length may vary considerably along a depth profile

(borehole). It tends to be smaller near the surface and gen-

erally increases with depth.

Deep fluid flow systems are controlled by steep or verti-

cal fracture systems that develop as a result of increasing

load with depth. Thermal water circulation in continental

basement is typically topography driven and follows the

vertical fracture systems to the surface where it may ulti-

mately discharge as a hot spring. The permeability struc-

ture dictates the ascent velocity of thermal water and

determines the final discharge temperature.

Faults or fault systems are prominent features of the

crust. The hydraulic conductivity of faults can be lower

than, equal to, or higher than that of the fractured rock

matrix. Examples of low-permeability faults have been pre-

sented in Stober & Bucher (2005a) and Stober et al.

(1999).

Permeability changes with time and the representative

volume is also temporal and variable. The transient nature



of j cannot normally be recorded by well test methods.

The variation of the permeability structure of the upper

continental crust is related to neotectonic processes and

the chemical interaction of fluid with the rocks it comes in

contact to along the flow path. Reactive fluid flow causes a

number of chemical effects. Dissolution and precipitation

of minerals on the fracture walls are relevant for the per-

meability of the system. These reactions tend to be kineti-

cally slow at temperatures below 200°C, corresponding to

about 5 km depths from where well test T/H data are

available. Thus, indications of permeability changes in nor-

mal continental crust from that depth range are scarce and

mainly inferential (e.g., long-term temperature changes of

hot springs). An analysis and discussion of permeability–

time relationships related to coupled thermal–hydrological–

mechanical–chemical processes can be found in Rudqvist

(2014).

In the depth range 5–15 km brittle deformation and

Darcy fluid flow still dominates. Evidence for transient per-

meability and its variability with time can be studied at

outcrops from this depth range exposed at the present-day

erosion surface. The evidence indicates that hydrothermal

fluid–rock interaction tends to first increase permeability

after initial fracture formation and then later reduce perme-

ability by depositing solid reaction products on the fracture

surfaces until the fracture is completely sealed and impervi-

ous to fluid flow. The typical time dependence of the con-

ductivity of a single fracture is mostly related to the

variation of the aperture with time. Flow at a given instant

in time can be approximated by the cubic law for fluid

flow, where the flow rate per hydraulic head difference is

proportional to the cube of fracture aperture. The fracture

aperture is a function of time due to mechanical aperture

variations such as extension, compression, shearing, and

other deformational effects, in addition to the progressing

chemical reactions. The permeability of a representative

volume in the lower part of the upper crust above the duc-

tile transition zone and below the reach of deep boreholes,

at a given instant in time, comprises the integral conduc-

tive property of all fractures that contribute to flow.

Although each single fracture probably has its conductiv-

ity–time relationship, the permeability of the representative

volume may not necessarily vary considerably with time.

The time dependence of the permeability in this part of

the upper continental crust is not realistically accessible

from hydraulic data, geophysical data, or from numerical

models. This rather pessimistic view is well grounded in

the extreme complexity of the processes controlling disso-

lution–precipitation reactions in hydrothermal environ-

ments. One often neglected complexity is the effects of the

Earth tidal forces, which keep even deep fluids in continu-

ous motion despite the lack of obvious forces for flow such

as topography in continuous motion. The tidal pumping

of fluids into reactive microporosity increases reaction

kinetics of fluid–rock interaction in a complex manner that

is difficult to quantify and to study experimentally. The

effect of tidal fluid motion in the upper crust is a feature

that deserves more attention in the geoscience community.

ACKNOWLEDGEMENTS

We thank Steve Ingebritsen and Tom Gleeson for taking

the initiative for this exciting special volume of Geofluids

and for taking the burden for all the editorial work this

involves. We acknowledge the very constructive reviews of

all five reviewers, which helped to improve the manuscript

substantially. We again thank Steve Ingebritsen for out-

standing editorial work on our manuscript, which he

greatly improved. Of course, all remaining errors remain in

the responsibility of the authors.

REFERENCES

Allen DM, Grasby SE, Voormeij DA (2006) Determining the

circulation depth of thermal springs in the southern RockyMountain Trench, south-eastern Columbia, Canada using

geothermometery and borehole temperature logs. HydrogeologyJournal, 14, 159–72.

Alt-Epping P, Diamond LW, H€aring MO (2013) Prediction ofwater–rock interaction and porosity evolution in a granitoid-

hosted enhanced geothermal system, using constraints from the

5 km Basel-1 well. Applied Geochemistry, 38, 121–33.Ankit K, Nestler B, Selzer M, Reichardt M (2013) Phase-field studyof grain boundary tracking behavior in crack-seal microstructures.

Contributions to Mineralogy and Petrology, 166, 1709–23.Aquilina L, Genter A, Elsass P, Pribnow D (2000) Evolution offluid circulation in the Rhine graben: constraints form the

chemistry of present fluids. In: Hydrogeology of Cristalline Rocks(eds Stober I, Bucher K), pp. 177–203. Kluwer Academic

Publishers, Dordrecht.Armbruster T, Kohler T, Meisel T, N€agler TF (1996) The zeolite,

fluorite, quartz assemblage of the fissure at Gibelsbach, Fiesch

(Valais, Switzerland): crystal chemistry, REE patterns, and

genetic speculations. Schweizerische Mineralogische undPetrographische Mitteilungen, 76, 131–46.

B€achler D (2003) Coupled Thermal-Hydraulic-Chemical

Modelling at the Soultz-sous-Forets HDR Reservoir (France).

Diss. ETH Zurich No. 15044, 151.Bauer F (1987) Die Kristallinen Gesteine aus der

Bohrlochvertiefung Urach 3 und ihre fluiden Einschl€usse: Eine

Interpretation der hydrothermalen €Uberpr€agung anhand derFluid-Daten aus Einschlußmessungen. Dissertation atUniversit€at (T.H.) Fridericiana Karlsruhe, p. 118.

Bear J (1979) Hydraulics of Groundwater. McGraw Hill Book

Comp, New York, NY.Bendall B, Hogarth R, Holl H, McMahon A, Larking A, Reid P

(2014) Australian experience in EGS permeability enhancement

– a review of 3 case studies. Proceedings, Thirty-Ninth Workshopon Geothermal Reservoir Engineering, Stanford University,Stanford, p. 10.

Berckhemer H, Rauen A, Winter H, Kern H (1997) Petrophysical

properties of the 9-km-dep crustal section at KTB. Journal ofGeophysical Research, 102, 18337–61.

Borchardt R, Emmermann R (1993) Vein minerals in KTB rocks.

KTB Report, 2, 481–8.



Borchardt R, Zulauf G, Emmermann R, Hoefs J, Simon K (1990)

Abfolge und Bildungsbedingungen von Sekund€armineralien in

der KTB-Vorbohrung. KTB Report, 90–4, 76–88.Brantley SL (2004) Reaction kinetics of primary rock-forming

minerals under ambient conditions. In: Surface and GroundWater, Weathering, and Soils (eds Drever JI, Holland HD,

Turekian KK), pp. 73–117. Treatise on Geochemistry, Elsevier,Amsterdam.

Bredehoeft JD (1967) Response of well-aquifer systems to Earth

tides. Journal of Geophysical Research, 72, 3075–87.Brown ET, Hoek E (1978) Trends in relationships betweenmeasured in-situ stresses and depth. International Journal ofRock Mechanics and Mining Sciences & Geomechanics Abstracts,15, 211–5.

Browne PRL, Courtney SS, Wood CP (1989) Formation rates ofcalc-silicate minerals deposited inside drillhole casing,

Ngatamariki Geothermal Field, New Zealand. AmericanMineralogist, 74, 759–63.

Bucher K, Stober I (2010) Fluids in the upper continental crust.

Geofluids, 10, 241–53.Bucher K, Zhang L, Stober I (2009) A hot spring in granite of

the Western Tianshan, China. Applied Geochemistry, 24, 402–10.

Bucher K, Stober I, Seelig U (2012) Water deep inside the

mountains: Unique water samples from the Gotthard rail base

tunnel, Switzerland. Chemical Geology, 334, 240–53.Bucher-Nurminen K (1981) The formation of metasomatic

reaction veins in dolomitic marble roof pendants in the Bergell

intrusion (Province Sondrio, Northern Italy). American Journalof Science, 281, 1197–222.

Bucher-Nurminen K (1989) Reaction veins in marbles formed by

a fracture-reaction-seal mechanism. European Journal ofMineralogy, 1, 701–14.

Butler JJ Jr (1998) The Design, Performance, and Analysis of SlugTests, p. 252. Lewis Publishers, New York, NY.

Caine JS, Evans JP, Forster CB (1996) Fault zone architecture

and permeability structure. Geology, 24, 1025–8.Cappa F, Rutqvist J, Yamamoto K (2009) Modeling crustal

deformation and rupture processes related to upwelling of deep

CO2-rich fluids during the 1965-1967 Matsushiro earthquakeswarm in Japan. Journal of Geophysical Research, 114, 1–20.

CFCFF (1996) Rock fractures and Fluid Flow, p. 551. National

Research Council, National Academy Press, Washington, DC.

Chebotarev II (1955) Metamorphism of natural waters in thecrust of weathering—1. Geochimica et Cosmochimica Acta, 8,22–48.

Connolly JAD, Podladchikov Y (2007) Decompaction weakening

and channeling instability in ductile porous media: implicationsfor asthenospheric melt segregation. Journal of GeophysicalResearch, 112, 1978–2012.

Cooper HH, Jacob CE (1946) A Generalized graphical method

for evaluating formation constants and summarizing well-fieldhistory. Transactions, American Geophysical Union, 27, 526–34.

Dietrich HG (1982) Geological results of the Urach 3 Borehole

and the Correlation with other Boreholes. In: The UrachGeothermal Project (ed. Haenel R), pp. 49–58. Schweizerbart’scheVerlagsbuchhandlung, Stuttgart.

Durney DW, Ramsay JG (1973) Incremental strains measured by

syntectonic crystal growths. In: Gravity and Tectonics (edsDeJong KA, Scholten R), pp. 67–96. Wiley, New York, NY.

Edmunds WM, Savage D (1991) Geochemical characteristics of

groundwater in granites and related crystalline rocks. In: AppliedGroundwater Hydrology, a British Perspective (eds Downing RA,Wilkinson WB), pp. 199–216. Clarendon Press, Oxford, U.K.

Elkhoury JE, Brodsky EE, Agnew DC (2006) Seismic waves

increase permeability. Nature, 441, 1135–8.Emter D, Wenzel HG, Z€urn W (1999) Das ObservatoriumSchiltach. DGG Mitteilungen, 3, 1–15.

Erzinger J, Stober I (2005) Introduction to Special Issue: long-

term fluid production in the KTB pilot hole, Germany.

Geofluids, 5, 1–7.Evans KF (2005) Permeability creation and damage due to

massive fluid injections into granite at 3.5 km at Soultz: 2.

Critical stress and fracture strength. Journal of GeophysicalResearch, 110, B04204.

Evans K, Wyatt F (1984) Water table effects on the measurement

of earth strain. Tectonophysics, 108, 323–37.Ferris JG (1951) Cyclic fluctuations of water level as a basis for

determining aquifer transmissivity. International Association ofHydrological Sciences (IAHS), 33, 148–55.

Filgris MN (2001) R€omische Baderuine Badenweiler. Historische

Wurzeln des Kurortes neu pr€asentiert. Denkmalsplege in Baden-W€urttemberg, 4, 166–75.

Forster C, Smith L (1989) The influence of groundwater flow on

thermal regimes in mountainous terrain: a model study. Journalof Geophysical Research, 94, 9439–51.

Frape SK, Fritz P (1987) Geochemical trends for groundwaters

from the Canadian shield. In: Saline Water and Gases inCrystalline Rocks (eds Fritz P, Frape SK), pp. 19–38. GeologicalAssociation of Canada Special Paper 33, The Runge PressLimited, Ottawa, ON.

Frost BR, Bucher K (1994) Is water responsible for geophysical

anomalies in the deep continental crust? A petrologicalperspective. Tectonophysics, 231, 293–309.

Garrels RM, Howard PF (1959) Reactions of feldspar and mica

with water at low temperature and pressure. In: Clays andClay Minerals (ed. Swineford AE.), pp. 68–88. InternationalSeries of Monographs on Earth Sciences, Pergamon Press,

Oxford, UK.

Genter A, Evans K, Cuenot N, Fritsch D (2010) Contribution of

the exploration of deep crystalline fractured reservoir of Soultzto the knowledge of enhanced geothermal systems (EGS).

Comptes Rendus Geoscience, 342, 502–16.H€afner F, Lauterbach M, Bamberg HF (1985) PhysikalischeEigenschaften von Grund- und mineralisiertem Wasser, von Erd€ol

und Erdgas. Wissenschaftlich-Technischer Informationsdienst, 26,18–27.

Hayba DO, Ingebritsen SE (1994) The computer modelhydrotherm, A three-dimensional finite-difference model to

simulate ground-water flow and heat transport in the

temperature range of 0 to 1,200°C, U.S. Geological Survey.

Water-Resources Investigations Report, 234, 94–4045.Horne RN (1996) Modern Well Test Analysis, A computer AidedApproach, 2nd edn, p. 250. Petroway Inc., Palo Alto, CA.

Hsieh PA, Bredehoeft JD, Rojstaczer SA (1988) Response of well

aquifer systems to Earth tides: problem revisited. WaterResources Research, 24, 468–72.

Ingebritsen SE, Manning CE (1999) Geological implications of a

permeability-depth curve for the continental crust. Geology, 27,1107–10.

Ingebritsen SE, Manning CE (2010) Permeability of the

continental crust: dynamic variations inferred from seismicity

and metamorphism. Geofluids, 10, 193–205.Iwatsuki T, Yoshida H (1999) Groundwater chemistry and

fracture mineralogy in the basement granitc rock in the Tono

uranium mine area, Gifu Prefecture, Japan-Groundwater

composition, Eh evolution analysis by fracture filling minerals.Geochemical Journal, 33, 19–32.



Jakob A, Mazurek M, Heer W (2003) Solute transport in

crystalline rocks at €Asp€o–II: blind predictions, inverse modelling

and lessons learnt from test STT1. Journal of ContaminantHydrology, 6, 175–90.

Kestin J, Khalifa HE, Correia RJ (1981) Tables of the dynamic

and kinematic viscosity of aqueous KCl solutions in the

temperature range 25–150°C and the pressure range 0.1–35 MPa. Journal of Physical and Chemical Reference Data, 10,57–70.

Kitagawa Y, Fujimori K, Koizumi N (2007) Temporal change in

permeability of the Nojima fault zone by repeated waterinjection experiments. Tectonophysics, 443, 183–92.

Konzuk JS, Kueper BH (2004) Evaluation of cubic law based

models describing single-phase flow through a rough-walled

fracture. Water Resources Research, 40, W02402, 1–17.Krusemann GP, de Ridder NA (1991) Analysis and evaluation of

pumping test data. International Institute for Land Reclamationand Improvement, 47, 377.

Kukkonen IK (1995) Thermal aspects of groundwater circulation

in bedrock and its effect on crustal geothermal modelling in

Finland, the central Fennoscandian Shield. Tectonophysics, 244,119–36.

Lee S-G, Kim T-K, Lee TJ (2011) Strontium isotope

geochemistry and its geochemical implication from hot spring

waters in South Korea. Journal of Volcanology and GeothermalResearch, 208, 12–22.

Lindberg A, Siitari-Kauppi M (1998) Shear zone-related

hydrothermal alteration in Proterozoic rocks in Finland. Water-Rock Interaction, WRI-9, 413–6.

L�opez DL, Smith L (1995) Fluid flow in fault zones: analysis of

the interplay of convective circulation and topographically driven

groundwater flow. Water Resources Research, 31, 1489–503.Markl G, Bucher K (1998) Composition of fluids in the lowercrust inferred from metamorphic salt in lower crustal rocks.

Nature, 391, 781–3.Mazurek M (2000) Geological and hydraulic properties of water-

conducting features in crystalline rocks. In: Hydrogeology ofCrystalline Rocks (eds Stober I, Bucher K), pp. 3–26. Kluwer

Acad. Publ., Dordrecht.

Mazurek M, Jakob A, Bossart P (2003) Solute transport incrystalline rocks at Aspo — I: geological basis and model

calibration. Journal of Contaminant Hydrology, 61, 157–74.Moore DE, Morrow CA, Byerlee JD (1983) Chemical reactions

accompanying fluid flow through granite held in a temperaturegradient. Geochimica et Cosmochimica Acta, 47, 445–53.

Nagra (1985) Sondierbohrung B€ottstein, Untersuchungsbericht.

Nagra Technischer Bericht (Technical Report of the Swiss NationalCooperative for the Disposal of Radioactive Waste), 85-01, p. 190.

Nagra (1992) Sondierbohrung Schafisheim, Untersuchungsbericht.

Nagra Technischer Bericht (Technical Report of the Swiss NationalCooperative for the Disposal of Radioactive Waste), 88-11, p. 193.

Nielsen KA (2007) Fractured Aquifers: Formation Evaluation byWell Testing, p. 229.Trafford Publishing, Victoria, BC.

Nishimoto S, Yoshida H (2010) Hydrothermal alteration of deep

fractured granite: effects of dissolution and precipitation. Lithos,115, 153–62.

Peters EJ (2012) Advanced Petrophysics: Volume 1: Geology,Porosity, Absolute Permeability, Heterogeneity, and Geostatistics,p. 238. Live Oak Book Company, Palo Alto, CA.

Polak A, Elsworth D, Yasuhara A, Grader AS, Halleck PM (2003)

Permeability reduction of a natural fracture under net dissolution

by hydrothermal fluids. Geophysical Research Letters, 30, 1–4.Priyatkina N, Kullerud K, Bergh S, Armitage P, Ravna E (2011)CO2 sequestration during interactions between fluid and

mafic to intermediate intrusive rocks on Vannøya Island,

West Troms Basement Complex, North Norway. Proceedings:10th International Congress of Applied Mineralogy, Trondheim(ICAM), pp. 549–53.

Ramsay JG (1980) The crack-seal mechanism of rock deformation.

Nature, 284, 135–9.Rogers PSZ, Pitzer KS (1982) Volumetric properties of aqueoussodium chloride solutions. Journal of Physical and ChemicalReference Data, 11, 15–81.

Rutqvist J (2014) Fractured rock stress-permeability relationships

from in situ data and effects of temperature and chemical –mechanical couplings. Geofluids, 15, 48–66.

Sanner B (2000) Baden-Baden, a famous thermal spa with a long

history. Geo-Heat Center Quarterly Bulletin, 21, 16–22.Stenger R (1982) Petrology and geochemistry of the basementrocks of the research drilling Projekt Urach 3. In: The UrachGeothermal Project (ed. Haenel R), pp. 41–8. Schweizerbart’scheVerlagsbuchhandlung, Stuttgart.

Stober I (1986) Str€omungsverhalten in Festgesteinsaquiferen mit

Hilfe von Pump- und Injektionsversuchen. Geologisches Jahrbuch,Reihe C, 42, 204.

Stober I (1995) Die Wasserf€uhrung des kristallinen Grundgebirges,p. 191.Ferdinand Enke Verlag, Stuttgart.

Stober I (1996) Researchers Study Conductivity of Crystalline

Rock in Proposed Radioactive Waste Site. EOS, Transactions ofthe American Geophysical Union, 77, 93–4.

Stober I (2011) Depth- and pressure-dependent permeability in

the upper continental crust: data from the Urach 3 geothermal

borehole, southwest Germany. Hydrogeology Journal, 19, 685–99.

Stober I, Bucher K (2004) Fluid Sinks within the Earth’s Crust.

Geofluids, 4, 143–51.Stober I, Bucher K (2005a) The upper continental crust, anaquifer and its fluid: hydraulic and chemical data from 4 km

depth in fractured crystalline basement rocks at the KTB test

site. Geofluids, 5, 8–19.Stober I, Bucher K (2005b) Deep-fluids: Neptune meets Pluto.In: The Future of Hydrogeology (ed. Voss C), pp. 112–5.Hydrogeology Journal.

Stober I, Bucher K (2007a) Hydraulic properties of the crystallinebasement. Hydrogeology Journal, 15, 213–24.

Stober I, Bucher K (2007b) Erratum to: hydraulic properties of

the crystalline basement. Hydrogeology Journal, 15, 1643.Stober I, Richter A, Brost E, Bucher K (1999) The OhlsbachPlume: natural release of deep saline water from the crystalline

basement of the black forest. Hydrogeology Journal, 7, 273–83.Sun H, Feistel R, Koch M, Markoe A (2008) New equations for

density, entropy, heat capacity and potential temperature of asaline thermal fluid. Deep-Sea Research. Part I, OceanographicResearch Papers, 55, 1304–1310.

T�oth J (1962) A theory of groundwater motion in small drainage

basins in central Alberta, Canada. Journal of GeophysicalResearch, 67, 4375–88.

T�oth J (1978) Gravity-induced cross-formational flow of

formation fluids, red earth region, Alberta, Canada: analysis,patterns, and evolution. Water Resources Research, 14, 805–43.

Wagner W, Kretschmar H-J (2008) International Steam Tables,Properties of Water and Steam, 2nd edn, p. 392. Springer-

Verlag, Berlin, Heidelberg.Weisenberger T, Bucher K (2010) Zeolites in fissures of granites

and gneisses of the Central Alps. Journal of MetamorphicGeology, 28, 825–47.

Weisenberger T, Bucher K (2011) Mass transfer and porosityevolution during low temperature water–rock interaction in



gneisses of the Simano nappe: Arvigo, Val Calanca, Swiss Alps.

Contributions to Mineralogy and Petrology, 162, 61–81.Xue L, Cast T (2013) Continuous permeability measurementsrecord healing inside the Wenchuan earthquake fault zone.

Science, 340, 1555–9.Yadav SK, Chakrapani GJ (2006) Dissolution kinetics of rock-water

interactions and its implications. Current Science, 90, 932–7.Yamamoto K, Yoshida H, Akagawa F, Nishimoto S, Metcalfe R

(2013) Redox front penetration in the fractured Toki Granite,

central Japan: an analogue for redox buffering in fractured

crystalline host rocks for repositories of long-lived radioactive

waste. Applied Geochemistry, 35, 75–87.Yardley BWD, Bodnar RJ (2014) Fluids in the continental crust.

Geochemical Perspectives, 3, 127.Yoshida H, Metcalfe R, Ishibashi M, Minami M (2013) Long-

term stability of fracture systems and their behaviour as flowpaths in uplifting granitic rocks from the Japanese orogenic

field. Geofluids, 13, 45–55.



Documents

Hydraulic conductivity of fractured upper crust: insights from hydraulic tests in boreholes and fluid-rock interaction in crystalline basement rocks